Atmos. Chem. Phys., 10, 1555–1576, 2010
www.atmos-chem-phys.net/10/1555/2010/
© Author(s) 2010. This work is distributed under
the Creative Commons Attribution 3.0 License.
Atmospheric
Chemistry
and Physics
The chemistry of OH and HO2 radicals in the boundary layer
over the tropical Atlantic Ocean
L. K. Whalley1,5 , K. L. Furneaux1,† , A. Goddard1 , J. D. Lee2,6 , A. Mahajan1 , H. Oetjen1 , K. A. Read2,6 , N. Kaaden3 ,
L. J. Carpenter2 , A. C. Lewis2,6 , J. M. C. Plane1 , E. S. Saltzman4 , A. Wiedensohler3 , and D. E. Heard1,5
1 School
of Chemistry, University of Leeds, Woodhouse Lane, Leeds, LS2 9JT, UK
Department, University of York, Heslington, YO10 5DD, UK
3 Physics Department, Leibniz Institute for Tropospheric Research, Leipzig, Germany
4 Department of Earth System Science, University of California, Irvine, CA, USA
5 National Centre for Atmospheric Sciences, University of Leeds, Leeds, LS2 9JT, UK
6 National Centre for Atmospheric Sciences, University of York, Heslington, YO10 5DD, UK
† Sadly passed away 28 July 2009
2 Chemistry
Received: 9 July 2009 – Published in Atmos. Chem. Phys. Discuss.: 28 July 2009
Revised: 5 January 2010 – Accepted: 19 January 2010 – Published: 15 February 2010
Abstract. Fluorescence Assay by Gas Expansion (FAGE)
has been used to detect ambient levels of OH and HO2 radicals at the Cape Verde Atmospheric Observatory, located
in the tropical Atlantic marine boundary layer, during May
and June 2007. Midday radical concentrations were high,
with maximum concentrations of 9 ×106 molecule cm−3 and
6×108 molecule cm−3 observed for OH and HO2 , respectively. A box model incorporating the detailed Master Chemical Mechanism, extended to include halogen chemistry, heterogeneous loss processes and constrained by all available
measurements including halogen and nitrogen oxides, has
been used to assess the chemical and physical parameters
controlling the radical chemistry. The model was able to
reproduce the daytime radical concentrations to within the
1 σ measurement uncertainty of 20% during the latter half
of the measurement period but significantly under-predicted
[HO2 ] by 39% during the first half of the project. Sensitivity analyses demonstrate that elevated [HCHO] (∼2 ppbv)
on specific days during the early part of the project, which
were much greater than the mean [HCHO] (328 pptv) used
to constrain the model, could account for a large portion of
the discrepancy between modelled and measured [HO2 ] at
this time. IO and BrO, although present only at a few pptv,
constituted ∼19% of the instantaneous sinks for HO2 , whilst
aerosol uptake and surface deposition to the ocean accounted
for a further 23% of the HO2 loss at noon. Photolysis of HOI
and HOBr accounted for ∼13% of the instantaneous OH formation. Taking into account that halogen oxides increase the
oxidation of NOx (NO → NO2 ), and in turn reduce the rate
of formation of OH from the reaction of HO2 with NO, OH
concentrations were estimated to be 9% higher overall due
to the presence of halogens. The increase in modelled OH
from halogen chemistry gives an estimated 9% shorter lifetime for methane in this region, and the inclusion of halogen
chemistry is necessary to model the observed daily cycle of
O3 destruction that is observed at the surface. Due to surface
losses, we hypothesise that HO2 concentrations increase with
height and therefore contribute a larger fraction of the O3 destruction than at the surface.
1
Introduction
The OH radical is one of the most dominant oxidants in the
troposphere. It is most prevalent in tropical regions where
high levels of humidity and solar irradiance lead to high rates
of OH production:
O3 + hv(λ < 340 nm) → O2 + O(1 D)
(R1)
O(1 D) + H2 O → 2 OH
(R2)
Correspondence to: L. K. Whalley
([email protected])
Published by Copernicus Publications on behalf of the European Geosciences Union.
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Aerosol
Loss
HOX
hv
XO
NO
O3
hv, H2O
OH
RH, O2
RO2
O3
NO, O2
O
Peroxides
HO 2/R 2
HO2 Hete
pro roge
ces neo
ses
us
Loss
CO,
O
2
Fig. 1. Chemical reaction scheme showing the reactions affecting OH and HO2 concentrations in the remote MBL. X=I, Br and R= alkyl
group.
Around 80% of methane, the third most abundant greenhouse
gas, is processed within the tropical troposphere by the OH
radical with the tropical marine lower troposphere contributing ∼25% to the total tropospheric methane removal rate
(Bloss et al., 2005a). The loss of methane is disproportionately high in the lower tropical troposphere compared to at
higher altitudes in the tropical troposphere owing to the large
temperature dependence of the reaction between methane
and OH; the reaction rate is ∼24 times greater at the surface
than at the tropopause in the tropics.
Despite the importance of the lower marine tropical troposphere, OH measurements in this region, and also those
of HO2 , which is closely coupled to OH through the reaction cycles shown in Fig. 1, are sparse. Aircraft measurements of OH and HO2 were made during the Pacific Exploratory Missions (PEM) (Hoell et al., 1996, 1997, 1999;
Raper et al., 2001), with OH midday concentrations between
6–8×106 molecule cm−3 recorded in the tropical marine
boundary layer (MBL) using the Selected Ion Chemical Ionisation Mass Spectrometer (SICIMS) technique (Mauldin III
et al., 1999, 2001). Modelling studies (using a photochemical box model and a photostationary state point model) indicated that Reaction (R2) was the major source of HOx
(OH+HO2 ) (contributing 81%), while the major sinks involved the formation and subsequent loss of the peroxides
of H2 O2 and CH3 OOH (Chen et al., 2001). The models under-predicted the measured MBL OH by 15–20%; no
halogen chemistry was included in these models. The primary focus of many of the PEM campaigns was the freetropopshere, and so measurements in the boundary layer
were limited. OH measurements using Differential Optical Absorption Spectroscopy (DOAS) have been made onboard the R/V Polarstern, as part of the Air Chemistry and
Lidar Studies of Tropospheric and Stratospheric Species on
Atmos. Chem. Phys., 10, 1555–1576, 2010
the Atlantic Ocean (ALBATROSS) project held in October/November 1996 (Brauers et al., 2001). OH data were collected over the latitudinal range 5◦ N to 40◦ S in the Atlantic.
OH displayed a distinct diurnal profile with a maximum
around local noon of ∼7×106 molecule cm−3 . A simple
photochemical box model, containing just CO and methane
as sinks for OH, on average underpredicted OH by 16%, but
was within the combined calibration errors associated with
the OH, CO and photolysis frequency measurements.
In background marine airmasses, which are low in volatile
organic compounds (VOCs) and NOx , with HOx generated
mainly by Reactions (R1)–(R2) and removed primarily by
the formation and subsequent loss of peroxides, modelled
and measured OH concentrations are in reasonable agreement. In contrast to OH, however, models often significantly overpredict HO2 in such airmasses when only this basic chemistry is considered (Sommariva et al., 2004, 2006;
Kanaya et al., 2007).
The Southern Ocean Photochemistry Experiment
(SOAPEX-2) was held at Cape Grim, in North Western Tasmania, Australia, in 1999 under extremely clean conditions
([NO]<3 pptv). Two chemical schemes were constructed
to describe the free-radical budget; one containing just the
oxidation of CO and methane, the other also describing
the oxidation of non-methane hydrocarbons (NMHC) that
were also measured (Sommariva et al., 2004). Box model
calculations of [OH] using these two mechanisms agreed
to within 5–10%, implying a minor role for NMHC as
an OH sink, and both over-predicted [OH] by 10–20%.
[HO2 ], however, was overestimated by up to 40%, but
model agreement was improved by assuming an HO2 uptake
coefficient to aerosol of unity (Sommariva et al., 2004).
This, however, is probably an unreasonable assumption, as
recent experimentally determined uptake coefficients for
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
HO2 on realistic aerosol surfaces range only from 0.01–0.1
(Thornton et al., 2008; Taketani et al., 2008, 2009). During
the summer 2002 NAMBLEX project, which took place
at Mace Head on the West coast of Ireland, the role of
halogen oxides, specifically IO and BrO, in the chemistry
of the MBL was confirmed (Bloss et al., 2005b; Sommariva et al., 2006; Smith et al., 2006). IO and BrO react
quickly with HO2 , with rate coefficients at 25◦ C of 8.6 and
2.1×10−11 molecule−1 cm3 s−1 for IO and BrO, respectively
(IUPAC, 2006). The HOI and HOBr formed may then either
undergo photolysis to yield OH (and a halogen atom) or
be taken up onto an aerosol surface. Reaction of HOI and
HOBr with halide ions in acidified aerosol can form volatile
inter-halogen species (Vogt et al., 1996; Pechtl et al., 2007)
which photolyse in the gas phase allowing gaseous halogen
chemistry to continue. The inclusion of a halogen scheme
has been shown to improve the HO2 modelled to measured
agreement at a number of marine sites (Bloss et al., 2005b;
Smith et al., 2006; Sommariva et al., 2006; Kanaya et al.,
2007), and could account for the model discrepancy reported
during the SOAPEX-2 experiment if realistic HO2 uptake
coefficients were instead used (Haggerstone et al., 2005).
At coastal locations such as Mace Head, the primary
source of IO is the photolysis of molecular iodine (SaizLopez and Plane, 2004) and, to a smaller extent, iodocarbons
(Chameides and Davis, 1980; Carpenter et al., 1999), which
are released by macro-algae exposed at low tide. The iodine
atoms generated react rapidly with O3 to generate the IO radical. At coastal sites where macro-algae are exposed, IO follows a distinct diurnal profile, with concentrations peaking
with low tide during the daytime (Alicke et al., 1999; Whalley et al., 2007). The impact of halogens on tropospheric
oxidants in regions without tidally exposed macro-algae, for
example the open ocean, which covers ∼70% of Earth’s surface, was until recently, unconfirmed. A number of theoretical and observational studies had inferred indirectly the possible global impact of halogens (Dickerson et al., 1999; Vogt
et al., 1999; Galbally et al., 2000; von Glasow et al., 2002;
von Glasow et al., 2004), but observations in the remote
MBL were limited (Allan et al., 2000; Leser et al., 2003).
In 2006/2007 an 8 month measurement series of IO and BrO
at the Cape Verde Atmospheric Observatory (CVAO) was reported (Read et al., 2008), exhibiting broad diurnal profiles
following light intensity; with typical peak concentrations of
∼1.4 and 2.5 pptv, respectively. CVAO is thought to be representative of the open ocean tropical MBL; ship and aircraft
measurements confirm that the measurements at the site are
representative of the surrounding area (Read et al., 2008; Lee
et al., 2009a) and hence these observations suggest a potentially global presence of halogen oxides. Although IO and
BrO are present at much lower concentrations than observed
at coastal sites, for example Mace Head ([BrO]max =6.5 pptv;
Saiz-Lopez et al., 2004, and [IO]max ∼30 pptv; Commane,
2009) and Roscoff, France ( [IO]max ∼28 pptv; Furneaux et
al., 2009), the inclusion of bromine and iodine chemistry was
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1557
necessary to model adequately the daily O3 destruction that
is observed at Cape Verde (Read et al., 2008).
For [NO]<∼20 pptv, O3 , an important greenhouse gas in
the troposphere, is photochemically destroyed (Penkett et al.,
1997; Lee et al., 2009b) (e.g. via Reactions (R1)–(R4) and as
shown in Fig. 1).
O3 + OH → HO2 + O2
(R3)
O3 + HO2 → OH + 2O2
(R4)
When halogen oxides are present, further O3 destruction may
occur via the reaction of a halogen atom (X) with O3 and
subsequent reaction of XO with HO2 , NO, or XO:
X + O3 → XO + O2
(R5)
XO + HO2 → HOX + O2
(R6a)
HOX + hv → OH +X
(R6b)
XO + NO → X + NO2
(R7)
XO + IO → X + OIO
(R8)
XO + XO → 2X + O2
(R9)
Surface deposition of O3 may also contribute to its destruction in the BL. At higher levels of NO the regime switches to
a net O3 production via the following reactions:
HO2 + NO → OH + NO2
(R10)
RO2 + NO → RO + NO2
(R11)
NO2 + hv(λ < 420 nm) → NO + O
(R12)
O + O2 + M → O3 + M
(R13)
Read et al. (2008) reported a consistent daytime O3 destruction cycle at the CVAO with an annually averaged loss of
3.3±2.6 ppbv d−1 . In this paper, we present OH and HO2
measurements made at the observatory as part of the Reactive
Halogen in the Marine Boundary Layer (RHaMBLe) project
(Lee et al., 2009a). The contribution towards the overall O3
destruction from the reactions of OH and HO2 radicals with
O3 is determined directly. A detailed model is used to investigate the impact of the reactions of HO2 with BrO and
IO, and the subsequent recycling to OH via the photolysis
of HOBr and HOI. Of interest is the contribution of halogen chemistry to the rate of photochemical destruction of O3
and the budgets of OH and HO2 radicals. The significance
of halogen chemistry for the global oxidising capacity and
hence the removal rate of methane is also examined.
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Fig. 2. Schematic of the FAGE laboratory set-up during RHaMBLe (see text for details).
2
2.1
Experimental
Measurement of OH and HO2 radicals
Fluorescence Assay by Gas Expansion (FAGE) provides a
powerful method for the detection of tropospheric OH and
HO2 radicals (Heard and Pilling, 2003). The technique
utilises the strong A2 6 + (v0 =0)←X2 5i (v00 =0) Q1 (2) transition at 308 nm for OH laser excitation. HO2 can also be
detected by this method following initial chemical conversion to OH by adding NO and subsequent detection. Laserinduced fluorescence experiments of this type are conducted
at low pressures (∼133 Pa) to extend the fluorescence lifetime of the OH radical sufficiently to allow the fluorescence
to be discriminated from the more intense laser radiation.
This is achieved by temporal gating of the detector; a channel photo multiplier (CPM) (Perkin Elmer C943P) is used.
Switching the CPM off during the laser pulse also helps to
prevent detector saturation (Smith et al., 2006). OH fluorescence and scattered light (laser and solar) are recorded by a
gated photon counter (SRS SR400) during a 500 ns integration period (Gate A), commencing approximately 50 ns after
the end of the laser pulse (FWHM∼35 ns). A small fraction
of the total laser scattered light, along with solar scattered
light and detector dark counts constitute a background signal
to the OH fluorescence in gate A. The contribution of the solar scattered light and dark counts can be determined using a
second photon counting gate (Gate B), delayed by 50 µs from
the first, at a time when the fluorescence and laser pulse have
Atmos. Chem. Phys., 10, 1555–1576, 2010
subsided. Tuning the laser away from resonance with the
OH transition (online), to a wavelength where OH does not
absorb (offline), allows the signal contribution of the laser
scatter to be determined.
Ambient OH and HO2 measurements have been made using the University of Leeds ground-based FAGE instrument
since 1996 (Heard and Pilling, 2003). Details of the instrument have previously been reported (Smith et al., 2006), and
are only considered briefly here. The operating conditions
encountered at the CVAO were challenging and a modified
operating procedure was necessary. The instrument is housed
in an instrumented 20 ft sea container (Fig. 2), and comprises
a laser system, an OH reference cell for calibration of the
laser wavelength, fluorescence cells which are housed on the
roof of the sea container at a height of 3.5 m, and data acquisition and calibration systems.
During the RHaMBLe project, the instrument was located
at the CVAO (16.85◦ N, 24.87◦ W), at Calhau on a northeast
facing rocky foreshore on the island of São Vicente within
the Cape Verde archipelago. The island shores shelve steeply
and very little tidal variability is apparent near to the site.
A detailed map of Cape Verde and the layout of the site
can be found in Lee et al. (2009a). The site is adjacent to
the ocean, the University of Leeds ground-based FAGE instrument was approximately 50 m from the water’s edge and
the inlet was positioned approximately 3.5 m above sea level
(a.s.l.). A wavelength tuneable Nd:YAG pumped Ti:Sapphire
laser (Photonics Industries DS 20-532) was used to generate UV light at 308 nm. The Nd:YAG produces ∼9 W of
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
532 nm radiation at a pulse repetition rate of 5 kHz, and
is used to pump a Ti:Sapphire laser which generates up to
1.6 W of broadband near-IR radiation. Selecting a wavelength close to 924 nm, by changing the incident angle of
an intra-cavity diffraction grating, and frequency tripling, by
passing the light through two non-linear harmonic generation stages, generates the required radiation close to 308 nm
for OH excitation. The time taken for the Ti:Sapphire laser to
build up sufficient gain for laser output was found to change
significantly in this campaign. When optimally aligned, the
delay between the 532 nm pump pulse and the Ti:Sapphire
laser output can be as short as 200 ns, but as laser optics
become damaged or alignment shifts (possibly due to temperature changes) this delay can increase. The time between
the Ti:Sapphire laser pulse, t308 , and the time when the fluorescence photons are counted (beginning of gate A) is critical, and must be maintained constant to ensure that the instrument sensitivity and background laser scatter does not
change. A counter timer (Agilent Technologies, 225 MHz
universal counter) triggered at the same time (t0 ), as the laser,
and coupled to a photodiode (Hamamatsu, S6468 series),
which monitors the onset of the pulse of frequency tripled
308 nm light (t308 ), was implemented during this project to
measure the delay t308 −t0 . This system allowed any drift in
t308 −t0 to be directly monitored in real time and a correction
added to or subtracted from the time of the trigger controlling the start of the fluorescence photon counting gate A. The
correction was applied prior to the beginning of each data acquisition cycle, which lasted ∼450 s.
During this campaign a number of the optical coatings
within the laser became damaged as a result of the high intensity laser beam hitting dust or sea-salt particles that had
deposited on the optics. Although this problem occurs, to a
certain extent, in all environments, the effect was particularly
exacerbated at the CVAO, which is exposed to the elements
being close to the water’s edge and subject to wind-speeds
regularly in excess of 10 ms−1 . The resultant optical damage
led to a reduction in laser power at 308 nm from ∼40 mW
to <20 mW as the project progressed. The stability of the
delay time t308 − t0 was significantly worse than had been
observed previously (or since). In addition, tuning the laser
wavelength by 0.006 nm (the interval between the online and
offline positions) significantly changed t308 − t0 and hence
the position of gate A relative to the laser. As a result, the
laser power and hence the laser scattered light changed significantly between the online and offline positions. The timing drift precluded OH measurements during the first half
of the project, and the HO2 measurements were subject to
a larger uncertainty during this period. The timing problem
was eventually ameliorated by reducing the interval between
the online and offline wavelengths to 0.004 nm, for which the
change in t308 − t0 was much reduced.
Upon exiting the laser, the light was split by a dielectrically coated beam-splitter (CVI Optics Ltd). 80% was
directed to a fluorescence cell via a fibre launcher (Elliot
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GoldTM series), fibre optic (Elliot Scientific) and collimation
assembly (Oz Optics), and 19% to an OH reference cell, designed to aid the precise tuning of the laser wavelength to
the peak of the Q1 (2) OH transition. A large concentration
of OH is generated by the pyrolysis of H2 O vapour, using a
flow of humidified air over a heated filament. Light exiting
the reference cell was directed on a UV sensitive photodiode
(New Focus 2032) to determine the laser power for subsequent normalisation of the ambient LIF signal; a cut-off filter
was placed in front of the photodiode to prevent visible light
reaching it. The remaining UV light (1%) was directed onto a
fast photodiode (Hamamatsu, S6468 series), which was used
as input to the counter-timer to correct for any drift in the
laser timing.
The laser light entered the detection cell on the roof of
the container via a baffled side arm and intersected the ambient air stream, drawn through a 0.8 mm diameter nozzle
sat on a 25 mm turret, on an axis perpendicular to the fluorescence detection axis containing the CPM. The laser light
then exited the cell via a second baffled side arm and was reflected onto a second UV sensitive filtered photodiode (New
Focus 2032), the output of which was monitored to check
good laser alignment was being maintained through the cell.
Ambient temperature fluctuations can cause the voltage output from photodiodes to drift, this precludes the use of the
detection cell photodiode (located on the roof) to normalise
the LIF signal to laser power. High solar intensities were
experienced at the Cape Verde site (the solar zenith angle at
solar noon in June at the site was ∼5◦ ) leading to a large
amount of solar scattered light detected within the cell by
the CPM. A solar shade, consisting of a sheet of black aluminium (10 cm×10 cm) was placed 15 cm above the nozzle,
which reduced the solar counts that reached the detector by
an order of magnitude. Without the shade, solar counts at
solar noon could reach >200 counts s−1 , the random fluctuations of which are larger than the typical signal from OH,
making measurements extremely difficult. The windspeed at
the site was ∼10 ms−1 (and rarely dropped below 5 ms−1 ),
thus the average time an air parcel spent under the shade was
∼0.01 s, much shorter than the lifetime of either OH or HO2
radicals, and should not perturb their concentrations.
Normally, the University of Leeds ground-based FAGE instrument utilises two fluorescence cells for simultaneous OH
and HO2 measurements (Smith et al., 2006). Due to low laser
power, coupled with a CPM failure in the OH cell, however,
only one cell was operable during the project and was used
instead to make alternating sequential measurements of the
two radicals. This set-up has been successfully utilised on
a number of previous field projects, for example during the
recent CHABLIS project (Bloss et al., 2007). Signals from
the photon counter were integrated for 1 s periods. Online
data were acquired for 200 s and consisted of 100 s of fluorescence from OH, followed by 100 s of fluorescence from both
OH and converted HO2 (NO added to the cell). The offline
signal was then collected for a further period of 100 s, with
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
NO added after 50 s to measure any additional signal caused
by the addition of NO, for example from enhanced scattering
(there was no extra signal). The [OH] did not vary significantly during a single data acquisition period, and, therefore,
the signal due to HO2 could be determined by subtracting the
OH online signal from the HOx online signal. In total, one
data acquisition cycle (including the time taken to find the
peak of the OH transition) took ∼450 s.
Calibrations were performed every two to three days
throughout the campaign. A zero air generator (EcoPhysics
PAG 003) was coupled to a zero air trap (Megatech CE500KF-O-4R Gas Purifier) which reduced the levels of NOx ,
VOCs and CO to <1 pptv and following humidification
in a liquid water bubbler, this flow was directed down a
60 cm long quartz glass tube (internal diameter 22 cm) past a
Hg pen-ray lamp (LOT Oriel model 6035), under roughly
laminar flow conditions (Reynolds Number = 646). The
184.9 nm radiation emitted by the pen-ray lamp photolysed
H2 O vapour, generating OH and HO2 (in the presence of O2 )
in equal quantities. A few ppbv of O3 was generated from O2
photolysis and the subsequent O+O2 recombination reaction,
and this was used as a chemical actinometer for the 184.9 nm
flux. When calibrating for HO2 a 100 sccm flow of CO
(BOC, 1% CO in N2 ) was added to the humidified air flow
to rapidly convert all OH radicals to HO2 (∼2.5×10−3 s OH
lifetime). Approximately 5 slm of the central portion of the
flow within the tube was drawn into the detection cell, via
the 0.8 mm sampling nozzle, the excess flow (∼7 slm) was
directed towards an O3 analyser (Thermo Environmental Instruments – TEI – 42C) and Dewpoint Hygrometer (General
Eastern 1311DR sensor). Determining [H2 O] and [O3 ] enabled the radical concentration to be calculated using:
[OH] = [HO2 ] =
[O3 ][H2 O]σH2 O φHOx
[O2 ]σO2 φO3 P
(1)
where σH2 O and σO2 are the 184.9 nm absorption cross
sections of H2 O vapour and O2 , determined experimentally to be (7.1±0.2)×10−20 molecule−1 cm2 and
(1.37±0.14)×10−20 molecule−1 cm2 , respectively, for the
same pen-ray lamp used during the field measurements.
φHOx =1 is the quantum yield for production of OH or HO2 in
the absence of added CO, and φHOx =2 for HO2 in the presence of added CO. P is the profile factor, which is discussed
further below. With the pen-ray lamp switched off ([HOx ]=0)
it was found that there was no additional signal above the
background when NO was added to the FAGE cell, ruling out
the presence of any significant artefact due to HOx production either directly from photolysis of impurities (e.g. HNO3 )
in the NO, or from the photolysis of species desorbed from
the walls after exposure to NO.
A correction needs to be applied to account for the
parabolic distribution of radial flow velocities across the calibration tube under laminar flow conditions. The central portion of the flow sampled by the detection cell has a higher
velocity and hence spends less time in the photolysis region
Atmos. Chem. Phys., 10, 1555–1576, 2010
compared with the remainder of the flow (excess) that is sampled by the O3 analyser from regions of the tube with a lower
velocity. This radial flow distribution leads to more O3 in the
excess air relative to that in the FAGE sampled air at the centre. Any variability in the flux of the 184.9 nm light across the
tube can also perturb the O3 and radical profile in the tube.
The profile factor, defined as P =[O3 ]excess /[O3 ]central , has
been experimentally measured as 1.57±0.09 (Smith, 2007),
and is used in Eq. (1) to calculate [OH] and [HO2 ].
The sensitivity of the cell towards OH and HO2 , in units
of cts s−1 mW−1 molecule−1 cm3 , is given by:
COH =
SHO2
SOH
and CHO2 =
[OH] × P wr
[HO2 ] × P wr
(2)
SOH and SHO2 are the measured fluorescence signals
(cts s−1 ) due to OH (in the absence of added CO and
NO) and HO2 (in the presence of added CO and NO),
respectively. Pwr is the laser power entering the cell
in mW, which decreased from ∼19 to ∼7 mW during
the RHaMBLe campaign. The average instrument sensitivity was 1.1×10−7 cts s−1 mW−1 [OH]−1 for OH and
4.4×10−8 cts s−1 mW−1 [HO2 ]−1 for HO2 , with the 2 σ uncertainty estimated as ∼40% for OH and HO2 (Smith, 2007).
The cell pressure was optimised for OH detection rather than
for conversion of HO2 . The detection limit (LOD) is given
by:
s
SNR
1
1 Sback
LOD =
(3)
+
P wr×C
m
n
t
where SNR is the signal-to-noise ratio (1), m is the number of online data points (100), n is the number of offline points (50), Sback is the background signal comprising laser scattered light (∼20 cts s−1 ), solar scattered
light (max. ∼20 cts s−1 at noon) and CPM dark counts
(0 cts s−1 ) and t is the integration of each data point
(1 s). Using the above values and Pwr = 9 mW, LODs of
1.1×106 molecule cm−3 and 2.8×106 molecule cm−3 were
calculated for OH and HO2 , respectively.
2.2
Ancillary measurements
A number of ancillary measurements were made to permit a
detailed characterisation of the atmospheric composition at
the site and to constrain the box model. Details of these
measurements are listed in Table 1 and are discussed further in Lee et al. (2009a). The majority of these measurements are part of a long term dataset made continuously at
the Cape Verde Observatory since October 2006. The halogen oxide measurements made using Long Path-Differential
Optical Absorption Spectroscopy (LP-DOAS) were part of
an 8 month dataset, covering the period November 2006 to
June 2007 (Read et al., 2008).
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Table 1. Listing of the concurrent measurements made during the RHaMBLe project and used to constrain the box model.
Measurements
Instrument
Typical noon
value
Averaging
Time
LOD
(2 σ )
Reference
O3
CO
CH4
NO
NO2
Ethane
TEI 42c, UV absorption
Aerolaser, VUV resonance fluorescence
Flask Samples, Gas Chromatography analysis
Chemiluminesence detector
Chemiluminesence detector
Gas Chromatography- Flame Ionisation Detection
(GC-FID)
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
GC-FID
LP-DOAS
LP-DOAS
LP-DOAS
35 ppbv
104 ppbv
1821 ppbv
2.3 pptv
11 pptv
845±26 pptv
1 min
1 min
–
1h
1h
1h
1 ppbv
–
–
1.5 pptv
4.1 pptv
2.5 pptv
(Lee et al., 2009a)
(Lee et al., 2009a)
(Lee et al., 2009a)
(Lee et al., 2009b)
(Lee et al., 2009b)
(Read et al., 2009)
50±1.7 pptv
2.2±0.2 pptv
3.9±0.3 pptv
119±3.3 pptv
21±1.4 pptv
22±0.9 pptv
10±1.0 pptv
826±137 pptv
749±115 pptv
487±45 pptv
1.3 pptv
2.9 pptv
328 pptv
1h
1h
1h
1h
1h
1h
1h
1h
1h
1h
20–30 min
20–30 min
20–30 min
2.5 pptv
2.5 pptv
2.5 pptv
2.5 pptv
2.5 pptv
2.5 pptv
1.0 pptv
18 pptv
7 pptv
3 pptv
0.3–0.5 pptv
0.5–1 pptv
200 pptv
Filter Radiometer
Scanning Mobility Particle Sizer (SMPS) &
Aerosol Particle Sizer (APS)
Campbell Met Station
3.4×10−5 s−1 a
1×10−6 cm2 cm−3
1 min
30 min
–
–
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2009)
(Read et al., 2008)
(Read et al., 2008)
(Mahajan et al.,
2010)
(Lee et al., 2009a)
(Allan et al., 2009)
76%
1 min
–
(Lee et al., 2009a)
Campbell Met Station
Campbell Met Station
296 K
1016 hPa
1 min
1 min
–
–
(Lee et al., 2009a)
(Lee et al., 2009a)
Propane
Iso-butane
n-butane
Acetylene
Ethene
Propene
Isoprene
Acetaldehyde
Methanol
Acetone
IO
BrO
HCHO
J(O1 D)
Aerosol
Relative
Humidity
Temperature
Pressure
2.3
Box model using the master chemical mechanism
A box model was used to calculate OH and HO2 concentrations, and contained a near explicit chemical scheme
for the oxidative degradation of C1–C5 hydrocarbons, extracted from the Master Chemical Mechanism (MCM)
version 3.1 (Saunders et al., 2003). The entire MCM
treats the degradation of 135 VOCs and considers oxidation by OH, O3 and NO3 . The degradation continues until CO2 and H2 O vapour form as the final oxidation products. Complete details of the kinetic and photochemical data used in the mechanism are available at the
MCM website (MCM, http://mcm.leeds.ac.uk/MCM/home).
A halogen chemical scheme (listed in supplementary information: http://www.atmos-chem-phys.net/10/1555/2010/
acp-10-1555-2010-supplement.pdf) was added to the MCM
to assess the effect of IO, BrO and related species on the
budgets of OH, HO2 and O3 concentrations. The mechanism
contains photolysis of O3 , NO2 , HOI, HOBr, IO, aldehydes,
ketones, H2 O2 and organic hydroperoxides (RO2 H), but only
measurements of j (O1 D) were used to constrain the model.
The other photolysis rates were calculated based on a twowww.atmos-chem-phys.net/10/1555/2010/
stream scattering model (Hough, 1988) and assuming clearsky conditions. Under cloudy conditions calculated J values were scaled to the ratio of observed j (O1 D) to clear sky
j (O1 D). Dry deposition terms for H2 O2 (1 cm s−1 ) (Junkermann and Stockwell, 1999), RO2 H (0.9 cm s−1 ) (Junkermann and Stockwell, 1999), HNO3 (0.8 cm s−1 ) (Ganzeveld
and Lelieveld, 1995), CH3 OH (0.09 cm s−1 ) (Carpenter et
al., 2004) and HCHO and aldehydes (0.33 cm s−1 ) (Brasseur
et al., 1998) were included in the model. The model was constrained to hourly measurements of a number of VOCs (listed
explicitly in Table 1), NOx , O3 , CO, CH4 , j (O1 D) and meteorological parameters. The midday concentrations/values
of these model constraints are given in Table 1 for information; there were, however, some deviations from these typical
midday values in the hourly measurements that were used as
model inputs. The halogen scheme was constrained using
measurements of the diurnal cycle of [IO] and [BrO] averaged for May 2007. It was necessary to use an average because of incomplete data coverage primarily arising from IO
and BrO being detected by DOAS in different wavelength
regions and, therefore, not being measured simultaneously.
A H2 concentration of 500 ppbv was used, similar to that
Atmos. Chem. Phys., 10, 1555–1576, 2010
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
measured in the remote MBL at other locations, for example
during SOAPEX-2 at Cape Grim (Sommariva et al., 2004),
and used in OH model calculations during the ALBATROSS
campaign (Brauers et al., 2001). A constant HCHO concentration of 328 pptv was assumed in the base-case model
scenario, based on average [HCHO] measured by LP-DOAS
during May at the observatory (Mahajan et al., 2010). Deviations from this background [HCHO] used (and the impact
upon HOx levels) are considered in Sect. 3.4. The unmeasured intermediate species, generated from the oxidation of
VOCs and other species used to constrain the model, were
initialised at zero. The simultaneous rate equations were
solved using the FACSIMILE integrator (Curtis and Sweetenham, 1987). The model was allowed to run to steady
state (∼4 days) to stabilise the concentrations of any unmeasured intermediate species, e.g. alkyl nitrates, unmeasured
carbonyls and peroxides, before the OH and HO2 modelled
data were compared to the measurements.
The rate of heterogeneous loss of HO2 and a
number of halogenated species (supplementary information: http://www.atmos-chem-phys.net/10/1555/2010/
acp-10-1555-2010-supplement.pdf) to aerosol surfaces
has been included in the model. A simple scheme was used,
based on a first-order loss to aerosol surfaces (Ravishankara,
1997):
cg Aγ
k 0 loss =
(4)
4
where cg is the mean molecular speed of the gas molecules
(cm s−1 ), calculated using Eq. (5):
s
8 RT
cg =
(5)
π Mw
k 0 loss is the heterogeneous loss rate, A is the aerosol surface area per unit volume, γ is the uptake coefficient, R is
the universal gas constant, T is temperature and Mw is the
molecular weight of the gas. More recently, this scheme
was extended by Sommariva et al. (2006) and Haggerstone et
al. (2005) to incorporate loss rates that were dependent upon
the measured aerosol size-distribution, using an expression
which incorporated the transition between control by gasphase diffusion and molecular uptake (Sander, 1999). In the
background marine environment, the particle diameters are
small. Under such conditions, mass transfer is limited by
interfacial mass transport and Eq. (4) provides a reasonable
estimation of the aerosol uptake coefficient (Haggerstone et
al., 2005).
The particle radius and number population, used to determine the aerosol surface area, were measured with a
Scanning Mobility Particle Sizer (SMPS) (0–1 µm) and a
Aerosol Particle Sizer (APS) (1–10 µm) (Allan et al., 2009)
at a height of 30 m. The average aerosol surface area
observed at the CVAO at this height during the project
was ∼1×10−6 cm2 cm−3 . For particles of 0–1 µm diameter (which encompasses the bulk of the aerosols at the site) a
Atmos. Chem. Phys., 10, 1555–1576, 2010
dry aerosol distribution was determined. At an average ambient humidity of ∼80% the radius of sea-salt aerosols are
expected to grow by a factor of ∼1.4 (Allan et al., 2009),
leading to an ambient aerosol surface area of roughly a factor of 2 (radius2 ) larger for this size range. A strong gradient in the aerosol surface area at the ground to a height of
∼12 m has recently been recorded at the site (von Glasow,
2009). It is estimated from these gradient measurements that
the aerosol surface area is approximately a factor of 2 greater
at the height of the FAGE nozzle (∼3.5 m) compared with
the aerosol surface area at 30 m. To account for the difference in sampling height and aerosol growth factors an aerosol
surface area of 4×10−6 cm2 cm−3 was used in model calculations. An uptake coefficient (γ ) of 0.1 was used based
on recent recommendations (Taketani et al., 2008). Sensitivity analyses have been performed on the aerosol uptake (see
Sect. 3.4 for further details).
A surface deposition term for HO2 was also considered.
The deposition velocity (VD ) of HO2 was estimated using
the NOAA/COARE air/sea gas exchange parameterization
(Fairall et al., 2007). The deposition velocity of HO2 is related to the flux (F ) of HO2 at a given height above the surface (z) by:
F = −VD × [HO2 ](z)
(6)
VD can be determined given knowledge of the sum of air-side
(Ra ) and water-side (Rw ) resistances (Fairall et al., 2000):
1
VD =
(7)
Ra + (Rw /α)
where α is the dimensionless solubility (a function of species,
temperature and salinity). The air-side resistance, Ra , represents the sum of the aerodynamic resistance (R1 ) and gasphase film resistance (R2 ). Assuming that the atmospheric
surface layer is neutrally stable then R1 can be determined
using Eq. (8):
U(z)
1
z
R1 = 2 =
ln
(8)
κu∗
z0
u∗
where U(z) is the mean wind-speed at height (z), κ is the
von Karman constant (0.4), u∗ , is the friction velocity, and
z0 , is the aerodynamic roughness length (Stull, 1988). The
gas-phase film resistance (R2 ) is determined by Eq. (9):
R2 =
5 2/3
S
u∗ C
(9)
where SC is the Schmidt number.
The water-side resistance (Rw ) can be determined by
Eq. (10) (assuming the influence of turbulence is negligible
relative to chemical reaction time-scale of HO2 in water):
1
Rw = √
aDw
(10)
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
1563
Fig. 3. Panel (A): Time-series of j (O1 D) (black line) and CO (blue line) from the HOx measurement period (gaps in data are due to power
failures). Panel (B): Time-series of NO (lime line), NO2 (green line) and O3 (red line) from the HOx measurement period (gaps in data are
due to power failures or instrument down-time). Panel (C): Time-series of measured HO2 in molecule cm−3 (pink line) with the 2 σ standard
deviation shown. Panel (D): Time-series of measured OH in molecule cm−3 (purple line) with the 2 σ standard deviation shown. All data
are hourly averaged.
where a is the reaction time-scale for HO2 in water (assumed
to be sufficiently fast that turbulence is neglected) and Dw is
the diffusion coefficient of HO2 in water.
Many of the variables in Eqs. (7)–Eq. (10) can be estimated by the NOAA/COARE algorithm given an input of
boundary layer depth (∼1000 m) (Read et al., 2008), atmospheric pressure (∼1010 mbar), wind-speed (0–20 m s−1 ),
height of wind-speed data (10 m), Henry’s law coefficient
(1.2×103 M atm−1 (Schwartz, 1984)) and rate coefficients
for reactions of HO2 in water (ionisation of HO2 form+
ing O−
2 and H ; Bielski, 1978, and chemical reaction of
−
HO2 with O2 forming HO−
2 ; Bielski, 1978, were considered). The water-side Schmidt number and diffusion coefficient were assumed to be similar to those of O3 (500 and
1.2×10−9 m2 s−1 ; Chang et al., 2004). The HO2 water-side
reaction rate was varied from 103 –105 s−1 to determine the
sensitivity of VD to it. For HO2 , the solubility is sufficiently
large that the deposition velocity is essentially air-side controlled and the flux is insensitive to the water side turbulence
and reaction rate.
3
Results
Figure 3 shows the values of j (O1 D), CO, NO, NO2 , O3 ,
HO2 and OH measured at the CVAO during the summer 2007
campaign. OH measurements were made on 5 days (28 May
to 2 June) and HO2 on 11 days (21 May to 2 June). From
21 May–24 May, the air-mass intercepted was characterised
www.atmos-chem-phys.net/10/1555/2010/
by elevated levels of CO (∼130 ppbv); CO levels dropped to
∼100 ppbv during the day on 24 May and remained at this
level for the remainder of the radical measurement period
(Fig. 3a). Back trajectory analysis for this period indicates
a switch in air-mass origin (Lee et al., 2009a). From 21–
24 May, the air intercepted had originated (at a low altitude)
over the western north Atlantic and continental America before being transported at a higher altitude over the ocean
and descending close to the Cape Verde islands. From 26–
29 May the air sampled originated over Southern Europe and
the Mediterranean (usually within the boundary layer), before passing near to the Canary Islands on its way to Cape
Verde. [CO] remained at background levels (∼100 ppbv)
between 26–29 May suggesting that this air-mass was not
strongly influenced by anthropogenic emissions from Europe. From 30 May onwards, the CVAO largely sampled
air that was mid north Atlantic in origin, passing close to the
Canary Islands before approaching Cape Verde from a North
Easterly direction. Throughout the measurement period the
maximum noontime j (O1 D) levels were relatively constant,
at 3.5×10−5 s−1 (Fig. 3a) with very little cloud cover observed throughout the campaign. O3 concentrations between
35–40 ppbv were recorded (Fig. 3b) (Lee et al., 2009a).
3.1
Daytime measured and modelled OH and HO2
During 3 out of the first 4 days of measurements [HO2 ]
was elevated relative to data taken later in the campaign,
corresponding to the elevated CO during this time, with a
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Fig. 4. Panel (A): Time-series of measured HO2 in molecule cm−3 (pink line) with the 2 σ standard deviation shown and MCM calculated
[HO2 ] (base-case scenario, see text for details) (black line). Panel (B): Time-series of measured OH in molecule cm−3 (purple line) with
the 2 σ standard deviation shown, and MCM [OH] calculated (base-case scenario, see text for details) (black line). All data represent a 1 h
average.
maximum of ∼6×108 molecule cm−3 (∼24.5 pptv) recorded
on 24 May (Fig. 3c); [NOx ] were elevated on this day and
may indicate a local pollution source. MCM model estimates of [HO2 ] over this period are also elevated relative to model estimates for later in the campaign (Fig. 4a),
although peak MCM modelled [HO2 ] from this period is
only ∼3×108 molecule cm−3 , suggesting either a missing
source of HO2 , or an over-estimation of the HO2 sinks in
the model; the modelled [HO2 ] estimates are ∼39% lower
than those measured during this phase (falling outside the
1 σ measurement uncertainty). The daily variability observed in [HO2 ] during this period is also not captured by the
model; on 22 May, measured [HO2 ] was significantly lower
at ∼3×108 molecule cm−3 , but this drop in [HO2 ] was not
reflected in the model. The lack of variability in the modelled
HO2 may be due, in part, to the use of a constant average diurnal cycle for IO and BrO, constant heterogeneous losses
and/or constant [HCHO]. Sensitivity analyses of [HO2 ] to
each of these modelled parameters have been performed and
are presented in Sect. 3.4.
After 26 May the peak measured [HO2 ] were lower ∼2–
4×108 molecule cm−3 , and are mirrored by the model. The
model is in much better agreement with [HO2 ] measurements
during this phase, overestimating [HO2 ] by 16% on average;
this is within the 1 σ measurement uncertainty of 20%.
OH measurements were also made during the second measurement period (Fig. 3c). Peak concentrations
of 6–9×106 molecule cm−3 were recorded; the midday
[OH] gradually decreasing over the 5 days. The model
Atmos. Chem. Phys., 10, 1555–1576, 2010
under-predicted [OH] on average by 18%, although again
this is within the 1 σ measurement uncertainty.
3.2
Nighttime measured and modelled HO2
Low concentrations (∼0.6 pptv) of HO2 were measured on
two nights (24 and 25 May); due to the instrumental difficulties, outlined in Sect. 2.1 no nightime OH measurements
were made. Model estimates of nighttime HO2 significantly
under-predict the observations, however (Fig. 5). As reported
by Read et al. (2008) a daily O3 destruction cycle is observable at the CVAO, with the majority of this loss due to photolysis of O3 and subsequent reaction of O(1 D) with H2 O
vapour and via catalytic cycles involving IO and BrO. The
O3 concentration is replenished during the night by entrainment of air from the free tropopshere that is richer in O3 ,
followed by advection to the site (Read et al., 2008). The
HO2 concentration was observed to decrease to a minimum
at ∼21:30 and then slowly increase throughout the remainder
of the night, analogous to the observed nighttime O3 profile,
suggesting that the entrained air during the night is providing
a source of radicals.
The nighttime generation of OH and HO2 in the MCM
box model is from reactions of O3 with alkenes, specifically propene (∼72%) and isoprene and its oxidation products (∼25%). The reaction of NO3 with alkenes is not a significant source of HO2 , unlike Mace Head (Salisbury et al.,
2001). The OH generated from O3 + alkene reactions reacts
with acetaldehyde and methane, leading to the formation of
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
1565
Fig. 5. Time-series of nightime [HO2 ] measured on two consecutive nights. The red line represents the measured HO2 , with the 2 σ standard
deviation shown. The grey and black lines represent modelled [HO2 ] when the model is left unconstrained to PAN and when [PAN] is held
at 100 pptv respectively.
peroxy radicals, which in turn, through self reaction or reaction with NO generate HO2 . Alternatively, OH reacts with
CO, generating HO2 directly. The OH and HO2 radicals are
strongly coupled at night with ∼37% of OH formed from
the O3 +HO2 reaction. The model underestimation of HO2 at
night may derive from an unmeasured alkene, possibly one
that is biogenic in origin. Phytoplankton in oceans are known
to produce a large suite of VOCs (Moore et al., 1994; Shaw
et al., 2003). It is likely, however, that this source would also
be present in the daytime, and so would have to be included
as an OH sink.
The entrainment of peroxyacetyl nitrate (PAN) may also
act as a nighttime (as well as daytime) source of radicals. In
the modelling presented thus far, PAN is unconstrained in the
model, and the modelled concentration peaks at ∼100 pptv
at ∼15:00, corresponding to the peak in NO2 and then decreases to ∼10 pptv during the night. The thermal decomposition lifetime of PAN is just 20 min at 300 K (Bridier et
al., 1991), and entrainment of air from the cooler free troposphere (where the PAN lifetime will be much longer) could
lead to the rapid decomposition of the molecule, leading to
an additional source of radicals. To test this hypothesis, the
PAN concentration was constrained in the model to a constant 100 pptv (during both the day and night). Running
the model with constrained PAN improves the modelled to
measured HO2 throughout the day and night (Fig. 5). Jacobi et al. (1999) have reported PAN concentrations measured during a cruise of the R/V Polarstern from Cape Town
(South Africa) to Bremerhaven (Gemany) in May/June 1998.
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Whilst cruising a maximum of 500 km west of the African
coast, corresponding to 20◦ W–15◦ E, (CVAO is ∼500 km
west of the coast) PAN concentrations north of 10◦ N (CVAO
is at 16◦ 510 49) only dropped below 100 pptv on two days.
In contrast to these findings, Gallagher et al. (1990) found
that between 7◦ N–64◦ N and 20◦ W–50◦ W in late summer
that PAN concentrations in the north Atlantic ranged from
<1 pptv–40 pptv. Similarly, Muller and Rudolph (1992) did
not measure PAN above the LOD of 10 pptv in the tropics
(30◦ N–30◦ S) when crusising at 30◦ W. The proximity of the
measurement sites to the African coast seems to strongly influence the PAN content of the sampled air-masses (Jacobi
et al., 1999). Trajectory analysis from the RHaMBLe campaign indicated that for 21–25 May (during which time the
nighttime measurements of HO2 were made) the air-masses
intercepted at the CVAO were influenced by the US continent. In such air-masses the PAN content may have been
enhanced, and, therefore, could act as an important radical
source. Measurements of PAN at the site are necessary to
confirm this hypothesis.
3.3
Rate of production and destruction analysis
Figure 6 shows the instantaneous rates of production (ROPA)
and rates of destruction analyses (RODA) for OH and HO2 at
solar noon (averaged for the period 12:00–13:00 LT – local
time) as determined by the MCM. The ROPA and RODA
are averaged over all days for which there are respective
measurements of OH and HO2 . Figure 7 shows the diurnal
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Fig. 6. Pie-charts showing the MCM (base-case scenario, see text for details) average diurnal modelled OH and HO2 sources and sinks
between 12:00–13:00.
variation of the MCM calculated rates of production and
destuction of OH and HO2 for different chemical processes,
again averaged over the days when there are radical measurements. It should be noted that several of the pathways serve
to inter-convert OH to HO2 , or vice versa, and so, overall, do
not lead to an increase or decrease in the total HOx .
The photolysis of O3 and subsequent reaction with H2 O
vapour represents the dominant OH source, accounting for
∼76% of its production at solar noon. The photolysis of
H2 O2 and CH3 OOH are much smaller OH sources (∼3%
combined); at steady state these peroxides reach levels of
∼500 pptv and ∼1 ppbv respectively in the model, similar in
magnitude to observations of peroxides at Cape Grim (Ayers
et al., 1996; Monks et al., 1998) and also in the Cape Verde
region during the ALBATROSS cruise (Weller et al., 2000).
The recycling of HO2 to OH, via reaction with NO and O3 ,
also act as small additional OH sources. The hypohalous
acids, HOI and HOBr, which are generated by the reaction of
HO2 with IO and BrO respectively (Fig. 1), may be photolysed to OH or taken up on aerosol surfaces. An uptake coefficient for HOI and HOBr loss onto aerosols of γHOX =0.061
was assumed, consistent with experimental data (Mossinger
and Cox, 2001). Under this scenario, the photolysis of HOI
and HOBr contributes ∼13% to the instantaneous OH formation. There is some uncertainty in the γHOX assumed, with
some studies speculating that it may be an order of magnitude higher in the MBL (Abbatt and Waschewsky, 1998;
Wachsmuth et al., 2002). The sensitivity of the modelled
[OH] to this parameter is investigated in Sect. 3.4.2.
Atmos. Chem. Phys., 10, 1555–1576, 2010
The reactions of OH with CO and acetaldehyde are the major OH losses (∼28% and ∼25% respectively). The source
of acetaldehyde at remote marine locations, such as Cape
Verde, has not been determined conclusively and remains
subject to debate. Acetaldehyde is formed by the oxidation
of primary hydrocarbons, it may derive from direct terrestrial emissions or from the photochemical degradation of organic matter in the oceans (Singh et al., 2001 and references
therein). Singh et al. (2003) have inferred from aircraft measurements of acetaldehyde that the surface waters of the Pacific are greatly supersaturated with acetaldehyde and may
act as a large oceanic source. Read et al. (2010) suggest that
an oceanic flux of acetaldehyde is a likely source of the observed acetaldehyde concentrations in the Cape Verde region
also. In contrast to acetaldehyde, the other oxygenated VOCs
that were measured (acetone and methanol) play only a minor role as an OH sink (contained within “other” in Fig. 6).
Other sinks include reaction with methane, H2 , O3 , HO2 ,
HCHO and peroxides. Reaction of OH with NO2 , C2 –C5
hydrocarbons (NMHC) and MCM calculated chemical intermediates from OH oxidation of NMHC (oxygenated VOCs
and organo-nitrates) all account for a small fraction of the
OH loss, and are contained within the “other” segment also.
HO2 sources are dominated by the reaction of OH
with CO (∼41%) and, to a smaller extent, CH3 O with
O2 (∼16%); the source of CH3 O is the reaction of
CH3 O2 with NO. The presence of a small quantity of
NO (∼1.5 pptv) at the site serves to slightly enhance HO2
levels overall, because although HO2 +NO removes HO2
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
1567
Fig. 7. Left: Diurnal variation of the rate of OH production and loss calculated using the MCM. Right: Diurnal variation of the rate of HO2
production and loss taken from the MCM (base case, see text for details).
(Reaction R10), this is offset at these low levels of [NO]
by a larger overall rate of RO2 +NO → RO+NO2 (followed by rapid HO2 production from RO+O2 ) leading to
a net production of HO2 (Reactions R11 and R14) (Poppe,
1999). If NO concentrations were to increase further,
[RO2 ] would decrease (Fleming et al., 2006) and HO2 +NO
(Reaction R10) would begin to dominate over RO2 +NO
(Reaction R11) (as k10 =8.9×10−12 molecule−1 cm3 s−1 is
>k11 =7.7×10−12 molecule−1 cm3 s−1 (IUPAC, 2006)). Under this regime, reactions involving NO would act as an HO2
sink rather than a source. The sensitivity of the radicals to
NOx levels is further discussed in the subsequent section
(Sect. 3.4.2).
RO + O2 → Aldehyde + HO2
(R14)
The reaction of HCHO with OH, combined with the photolysis of HCHO, can contribute close to 20% of the HO2 at solar
noon. Other smaller HO2 sources include OH+H2 , OH+O3 ,
CH3 OH+OH, j (CH3 CHO), OH+H2 O2 and HO2 NO2 thermolysis.
The reaction of HO2 with IO and BrO accounts for a
combined, instantaneous, HO2 loss of ∼19%, despite these
species only being present at ∼1.4 pptv and ∼2.5 pptv (peak
daytime levels) respectively. The major HO2 sink is the reaction with RO2 (R=CH3 –C5 H11 ) (Reaction R15), typical of
low NOx environments (Carpenter et al., 1997; Penkett et al.,
1997), which together with the self reaction (Reaction R16)
accounts for over 40% of the total HO2 loss.
HO2 + RO2 → RO2 H + O2
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(R15)
HO2 + HO2 → H2 O2 + O2
(R16)
Heterogeneous loss to aerosol and surface deposition accounts for ∼23% of the HO2 loss at noon-time.
Over a 24 h period, the relative contribution of the HOx
sources and sinks remains fairly consistent, as shown in
Fig. 7. There is some deviation from the trend discussed
above, and displayed in Fig. 6, for example the reaction of
OH with CH3 CHO and reactions involving HO2 NO2 peak
in the afternoon, reflecting the afternoon peak in the concentrations of CH3 CHO and NO2 as measured at the site. In
the morning and evening (when the concentration of HO2 is
low) the reaction of HO2 with O3 dominates over the HO2
self reaction promoting radical recycling rather than radical
termination. The reaction of HO2 with IO and BrO exhibits
a broad diurnal cycle that reflects light intensity.
3.4
Sensitivity analysis of modelled OH and HO2
concentrations
The mean diurnal profiles of the observed and modelled
[OH] and [HO2 ] has been calculated (Fig. 8) over all days
for which there were respective measurements of OH and
HO2 . Owing to instrumental difficulties outlined in Sect. 2
only daytime OH data are available. In general good agreement between the modelled and measured [OH] profile is
achieved (Fig. 8a and C). The model under-predicts [OH]
in the mid-morning, and during the very late afternoon and
evening, although for the latter, OH levels are close to the
LOD of the instrument. HO2 was measured on 11 days and
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Fig. 8. Panel (A): Average diurnal [OH] calculated for the 5 days of measurements (data on an hourly time-scale); the error bars represent
the standard deviation of the data that were averaged (purple points and line). The modelled average diurnal (averaged over the 5 days of OH
measurements) is also shown (MCM= grey line). Panel (B): Average diurnal [HO2 ] calculated for the 11 days of measurements (data on an
hourly time-scale); the error bars represent the standard deviation of the data that was averaged (pink points and line). The modelled average
diurnal (averaged over the 11 days of HO2 measurements) is also shown (MCM= black line). Panel (C): Solid line, with black squares
represents the (OH MCM modelled)/(OH measured) ratio averaged for 1 h over 24 h. Panel (D): Solid line represents the (HO2 MCM
modelled)/(HO2 measured ratio) averaged for 1 h, for a 24 h period. Red line in panels C and D represents modelled : measured ratio of 1.
2 nights, enabling a complete 24 h average diurnal profile to
be built up. Excellent agreement between observed and modelled [HO2 ] is achieved (Fig. 8b and d) during the daytime;
during the nightime the [HO2 ] is replicated when the model
is constrained to a constant [PAN] of 100 pptv (Sect. 3.2).
As highlighted in Figs. 3 and 4 (and discussed in Sect. 3.1),
significant variability was observed in the measured daily
peak [HO2 ], certainly when different air-masses were intercepted; this variability was not replicated in model simulations to the same extent and so the agreement between the
modelled and measured [HO2 ] in Fig. 8 may be slightly fortuitous. It should be noted, however, that owing to the more
extensive data coverage during the latter half of the measurement period, when the agreement between the daily modelled and observed [HO2 ] was at its best (Fig. 4), the diurnal profiles will be more heavily weighted to this phase.
Several key modelled parameters that control the production
and loss of HO2 have been identified in Sect. 3.3, a number of which were either held at constant values ([HCHO]
and heterogeneous chemistry) or used a constant diurnal profile ([XO]) within the model. A number of model scenarios in which these parameters are varied have been run to
Atmos. Chem. Phys., 10, 1555–1576, 2010
determine the impact of these on the modelled [HO2 ] (and
[OH]) and to assess whether varying these parameters could
account for the daily variability in the observations. Table 2
lists the changes made and their impact on modelled [HO2 ],
and Fig. 9 shows the impact on the diurnal variation of the
modelled-to-measured ratio of [HO2 ]. The uptake coefficient of the hypohalous acids has also been varied, in accord
with reported literature values, to determine the sensitivity of
[OH] (and [HO2 ]) to this:
3.4.1
HCHO
As discussed in Sect. 3.3, the reaction of OH with HCHO
and HCHO photolysis are sources of HO2 . The model thus
far has been constrained to a constant HCHO concentration
of 328 pptv, which is the mean HCHO concentration observed during May 2007, as measured using the LP-DOAS
during its 8 month deployment at the observatory. It should
be noted, however, that throughout the 8 months [HCHO]
were variable, dropping below the LOD of the LP-DOAS
(∼200 pptv) at times and elevated in polluted air-masses.
Recent analysis of the HCHO data during the RHaMBLe
measurement intensive has highlighted that [HCHO] were
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
1569
Table 2. List of parameters that have been varied to optimise the HO2 modelled to measured agreement. The daytime % change was averaged
from 08:44–19:44; the nighttime % change was averaged from 19:44–08:44.
Model Parameter
Parameter Change
Av. % Change
in HO 2 (day)
Av. % Change
in HO2 (night)
Av. % Change
in OH (day)
HCHO
HCHO
Halogen chemistry
Halogen chemistry (modified
[NOx ]
Hypohalous acid chemistry
Aerosol uptake
Surface deposition
328 pptv → 100 pptv
328 pptv →2000 pptv
XO chemistry switched off
XO chemistry switched off
& [NOx ] → doubled
γHOX =0.061 → 0.61
Included → Not included
Included → Not included
−8
+60
+12
+18
0
+1
+2
0
+2
−9
−12
−9
−2
+29
+9
0
+59
+16
−4
+7
+2
Heterogeneous losses
& halogen chemistry
No heterogeneous loss,
no IO chemistry, [BrO]×1.4
+52
+104
−3
Fig. 9. (HO2 MCM modelled)/(HO2 measured) ratio for a number of different model scenarios. Black line = base case scenario (with
[PAN]= 100 pptv) , all other scenarios are based on this model with the specified (see legend) parameters adjusted. Red line represents an
(HO2 MCM modelled)/(HO2 measured) of 1. See text and Table 2 for further details.
extremely elevated (∼2 ppbv) on the 23 and 24 May (and
also on the 27 and 28 May, although HO2 data is limited
during this second period due to power cuts). These large
excursions from the more typical [HCHO] used in the model
could account for the model’s failure to replicate the elevated
[HO2 ] observed at the beginning of the measurement period
(Fig. 4).
Two model runs whereby [HCHO] was reduced to
100 pptv (Fig. 9, grey line) and increased to 2 ppbv (Fig. 9,
pink line) have been performed. Reducing the HCHO
concentration from 328 pptv to 100 pptv decreases [HO2 ]
by ∼8% during the daytime, but does not effect the nighttime levels of HO2 (Table 2); at [HCHO] of 2 ppbv, [HO2 ]
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increases by ∼60% on average from the base-case scenario.
Approximately a 140% increase in the base-case model is
required to match the elevated [HO2 ] observed at the beginning of the measurement period. This remaining discrepancy may indicate that the model is missing other additional
HO2 sources at this time or potentially that the HO2 sinks
are over-estimated by the model at this time (see Sect. 3.4.2
and 3.4.3). During the ALBATROSS cruise, up to 1 ppbv of
HCHO was observed in the tropical mid-Atlantic (Brauers
et al., 2001) and ∼750 pptv was recorded in the Cape Verde
region (Weller et al., 2000). At Cape Grim, Tasmania during
the SOAPEX-2 campaign, during clean conditions, [HCHO]
was measured around noon from 217–352 pptv (with 50%
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
error) (Sommariva et al., 2004). The source of the high
[HCHO] observed during RHaMBLe may be caused by longrange transport of polluted air from America; this will be
further investigated in future publications (Mahajan et al.,
2010).
3.4.2
Halogen chemistry
The inclusion of a halogen scheme has been shown to improve the HO2 modelled to measured agreement at a number of marine locations (Bloss et al., 2005b; Smith et al.,
2006; Sommariva et al., 2006; Kanaya et al., 2007). Sommariva et al. (2006) report that combined, IO and BrO accounted for ∼30% of the HO2 loss term during the NAMBLEX campaign at [IO]=4 pptv and [BrO]=6.5 pptv. Switching the halogen chemistry off in the Cape Verde model serves
to increase [HO2 ] by ∼12% during the daytime and ∼2%
during the night. [OH] decreases by ∼12%. Sommariva et
al. (2006) found that the modelled to measured HO2 agreement could be improved further during the NAMBLEX campaign by making the assumption that the source of IO was not
evenly distributed over the DOAS light-path, rather it was
concentrated by the shore close to the HO2 FAGE measurements, and as a consequence [IO] was ∼10 times higher by
the FAGE inlet. There are no macro-algae beds around the
Cape Verde islands unlike Mace Head that could be acting as
a source of IO and so it is unlikely that there is a hotspot of
IO production by the CVAO; this is supported by the good
modelled to measured HO2 agreement that exists when it is
assumed that the [IO] is evenly distributed throughout the
DOAS lightpath, representative of an oceanic source. If there
were some variability in the oceanic emissions from day to
day this could account for some of the variability observed
in the HO2 concentrations.
As discussed in Sect. 3.3, the overall impact of halogen chemistry on the OH concentration will be dependent
upon the uptake coefficient assumed for the hypohalous acid
species. To investigate the sensitivity of [OH] and [HO2 ] to
this parameter, a model run was performed in which γHOX
was increased 10 fold from the base-case run (see Table 2).
The [OH] was found to decrease by 4% during the day and so
under this scenario the impact of halogen oxides upon [OH]
and on the oxidising capacity is reduced, although the photolysis of HOI and HOBr is still acting as a significant OH
source. The [HO2 ] during the daytime is also reduced by 2%
under this scenario as a consequence of the reduction in the
rate of reaction of OH with CO.
The presence of halogen oxides not only serves to modify
the HOx chemistry, but also the NOx chemistry by accelerating NOx oxidation through Reactions (R7) and (R17). Thus
far, the impact of IO and BrO on NOx chemistry has not been
considered, yet any change in the NO:NO2 ratio can perturb
the OH:HO2 ratio by changing the rate of Reactions (R10)
and (R11).
XO + NO2 + M → XONO2 + M
Atmos. Chem. Phys., 10, 1555–1576, 2010
(R17)
Keene et al. (2009) report a reduction in [OH] in a modelling
study when halogen oxide chemistry was included as a consequence in the reduction in the secondary formation of OH
via Reaction R10 (Keene et al., 2009). Modelling studies
indicate that NOx levels could approximately double in the
absence of halogen oxides in environments similar to Cape
Verde (Keene et al., 2009). To simulate the overall impact
of halogen oxides, taking into account that NOx levels may
be significantly perturbed due to the presence of halogens,
a model run, initialised with no halogens and a factor of
two increase in the NOx concentration, has been performed.
In contrast to Keene et al. (2009), OH concentrations were
found to be 9% larger in the presence of halogen oxides compared to in the absence of halogen oxides (but with double
[NOx ]) (Table 2); the perturbation of the NOx chemistry by
halogen oxides offsets the instantaneous OH formation via
HOI and HOBr photolysis by ∼3%. Keene et al. (2009) consider the impact of BrO and ClO only; no IO measurements
were made. The rate coefficient for reaction of ClO with
HO2 (5×10−12 molecule−1 cm3 s−1 ; IUPAC, 2006) is a factor of two times smaller than the rate coefficient for reaction
of ClO with NO (1.7×10−11 molecule−1 cm3 s−1 ; IUPAC,
2006). Depending on [HO2 ] and [NO], the presence of ClO
can effectively lead to a reduction in [OH], as presented by
Keene et al. (2009). Conversely, the rate coefficient for reaction of IO with HO2 (8.4×10−11 molecule−1 cm3 s−1 ; IUPAC, 2006) is larger than the rate coefficient for reaction
of IO with NO (1.96×10−11 molecule−1 cm3 s−1 ; IUPAC,
2006) and so the presence of IO, in low NOx environments,
serves to increase [OH]. The absence of IO chemistry in the
model used by Keene et al. (2009) could account for the differences in the modelled impacts of halogens on [OH]. The
difference between the two model scenerios in which halogen chemistry is switched off and NOx chemistry is (a) constrained to observations and (b) is doubled, allows the impact of variable [NOx ] upon [OH] and [HO2 ] to be assessed.
On average, modelled [OH] and [HO2 ] increased by a further 3% and 6% respectively via increases in Reactions (R10)
and (R11) when [NOx ] was set to twice the observations. As
discussed above, halogen oxides can perturb this chemistry
and need to be condsidered alongside NOx to predict how
potential changes in [NOx ] may impact [HOx ] in this environment.
3.4.3
Heterogeneous losses of HO2
Heterogeneous loss of HO2 acts as a large HO2 sink
(Sect. 3.3). If the uptake onto aerosol surfaces is assumed
negligible, modelled [HO2 ] increases by 29% during the
daytime and by 59% at night, leading to a large modelled
HO2 over-prediction. A change in the aerosol surface area
and/or composition could account for some of the daily variability observed in HO2 . Thornton et al. (2008) report that
the uptake coefficient for HO2 may vary considerably with
changes in aerosol composition. Certainly, if heterogenous
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
1571
Fig. 10. Pie-charts highlighting the key parameters that cause the daily O3 destruction that is observed at the CVAO both at the surface (left
panel) and at the top of the MBL (right panel). Total O3 destruction calculated using 12 h averaged, daytime, data was ∼4 ppbv at the surface
and ∼3.5 ppbv at the top of the MBL.
loss processes to aerosol surfaces were lower during the interception of the long-range transported airmass compared
with loss processes during the interception of tropical Atlantic marine air then this (combined with elevated [HCHO])
may account for the model under-prediction of [HO2 ] at the
beginning of the campaign. Furthermore, the sensitivity of
modelled [HO2 ] to aerosol uptake, coupled with the steep
gradient in aerosol surface area from the ground to a height
of 30 m (Sect. 2.3) suggests that [HO2 ] could vary significantly with altitude within the BL.
In addition to heterogeneous loss to aerosols, surface deposition of HO2 to the surrounding ocean also contributes to
the total HO2 sink within the model, albeit to a smaller extent, and may also promote variability in [HO2 ] with altitude.
Without a surface deposition term, modelled [HO2 ] is ∼9%
higher during the daytime and ∼16% at night. Prior to reaching the FAGE cell, the sampled air passed over ∼50 m of surf
zone and rocks, in addition to the open ocean, and as this
area is likely to have an increased surface roughness, leading
to higher turbulence and hence contact with the ground, the
surface deposition assuming loss to the ocean only is likely
a lower limit for this region. The uncertainties in the magnitude of the surface deposition term could affect the overall
magnitude of the modelled [HO2 ], but are unlikely to account
for the daily variablility in the measurements at the ground.
4
Discussion and implications for MBL photochemistry
Tropospheric O3 is an important greenhouse gas, providing a
similar radiative forcing to methane, and ∼25% of that due to
CO2 (Forster et al., 2007), and in the troposphere forms via
the catalytic oxidation of CO, methane and other VOCs in
the presence of NOx . O3 production predominantly occurs
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in continental regions where NOx is elevated, whereas in
the remote tropical marine boundary layer O3 is destroyed
primarily through photolysis and subsequent reaction of the
excited state oxygen atom with H2 O vapour (Read et al.,
2008). During the RHaMBLe project at the CVAO, a daily
O3 loss was observed of the order of ∼3.5 ppbv d−1 . Taking
the daytime average concentrations for the relevant measured
species from the period 21 May–2 June 2007, the relative
contributions to the observed chemical loss were calculated
(concentration × rate coefficient), as shown in Table 3 and
Fig. 10. The photolysis of O3 and subsequent reaction with
H2 O (Reactions R1 and R2) accounts for ∼40% of the daily
O3 destruction. The reaction of O3 with OH (9%) (Reaction R3) and HO2 (12%) (Reaction R4) also contributes to
its loss, owing to the relatively high concentrations of the
radicals at the site, and reaction with IO (22.5%) and BrO
(16.5%) also destroy O3 (Reactions R5–R9). Table 3 highlights the relative importance of the reactions which contribute to the chemical destruction of O3 . When combined,
BrO and IO together destroy ∼1.5 ppbv d−1 , whilst the loss
due to the combined reaction with OH and HO2 is just under
1 ppbv d−1 , in good agreement with estimates reported by
Read et al. (2008) for the month of May. Read et al. (2008),
however, used a box model to calculate OH and HO2 for the
purposes of calculating the rate of O3 destruction, rather than
using the measured values directly in the field, as presented
here.
Aircraft measurements of O3 made above the CVAO
during the RHaMBLe project (Read et al., 2008) demonstrated that the O3 destruction observed at ground level occurs throughout the entire MBL, which typically extends to
∼1 km altitude. The IO and BrO, observed by the LP-DOAS
at the site, originate from the photolysis of organo-iodine
compounds volatilised from the ocean and the acid-catalysed
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
Table 3. The rates of the rate-limiting O3 destruction steps and their percentage contribution towards O3 destruction both in the surface layer
and at the top of the MBL (1 km), see text for details.
O3 destruction
rate determining
step
O1 D+H2 O
OH+O3
HO2 +O3
IO+IO
IO+HO2
IO+NO
IO+BrO
BrO+BrO
BrO+HO2
BrO+NO
Total
Contribution to O3 destruction
at the surface
(molecule cm−3 s−1 )
Contribution to O3
destruction at the top of the
BL (molecule cm−3 s−1 )
9×105 (40%)
2.1×105 (9%)
2.6×105 (12%)
7.2×104 (3%)
3.6×105 (16%)
2.7×104 (1%)
1.1×105 (5%)
1.4×104 (1%)
2.1×105 (10%)
6.7×104 (3%)
2.2×106 (3.91 ppbv d−1 )
9×105 (45%)
2×105 (9%)
3.8×105 (18%)
0%
0%
0%
0%
2.7×104 (1%)
4.5×105 (22%)
9.4×104 (5%)
2.0×106 (3.59 ppbv d−1 )
activation of bromine from sea-salt aerosols, respectively
(Read et al., 2008). There is some evidence to suggest
that [BrO] may increase with altitude in the MBL (von
Glasow et al., 2002) owing to an increase in the acidity of
sea salt aerosols with altitude which enables acid catalysed
bromine activation to occur more efficiently (von Glasow
and Sander, 2001). It is uncertain, given the surface source
of organo-iodine compounds and the short atmospheric lifetime of IO (∼50 s), whether this molecule also prevails
throughout the MBL. The rate of heterogeneous loss of HO2
by uptake to aerosol surfaces and by surface deposition to
the ocean (and/or rocky foreshore) will be at a maximum
at ground level. The HO2 measurements were made just
3.5 m a.s.l. and, as discussed above, heterogeneous losses
could reduce [HO2 ] at the ground by ∼38% during the daytime. As a consequence of this, the HO2 concentration could
increase with altitude in the boundary layer, and hence the
rate of destruction of O3 by the HO2 +O3 reaction will also
increase with altitude. Further increases in [HO2 ] with height
may be expected as a consequence of a reduction in the [IO]
with height.
A model scenario in which [IO], and heterogenous losses
are set to zero and [BrO] is increased by a factor of 1.4
to simulate chemistry at the top of the MBL predicts that
[HO2 ] could be 52% higher than at the surface during the
daytime (Table 2). Table 3 and Fig. 10, right panel, highlight the impact these changing parameters may have upon
O3 destruction. A reduction in temperature (∼5◦ C) is also
considered as this has a small effect on the rate coefficients
used. An O3 loss rate of ∼3.91 ppbv d−1 is predicted at the
surface and an O3 loss rate of ∼3.59 ppbv d−1 at the top
of the MBL (∼1 km). Although these analyses only consider the chemical processes that lead to a loss in O3 and
neglect physical processes, such as surface deposition and
Atmos. Chem. Phys., 10, 1555–1576, 2010
entrainment of O3 , which can modify the overall O3 destruction, this work demonstrates the increase in HO2 and BrO
concentrations is sufficient to compensate for the loss of IO
at the higher altitudes. Aircraft measurements of HO2 made
over the Pacific during the Transport and Chemical Evolution over the Pacific (TRACE-P) campaign (Cantrell et al.,
2003) provides some evidence that HO2 concentrations increase over the first couple of kilometres. However, it is
difficult to draw strong conclusions from the dataset due to
the limited number of boundary layer observations that were
made. Leser et al. (2003) have reported MAX-DOAS measurements of BrO taken onboard the R/V Polarstern north of
the Canary Islands. Significant boundary layer BrO was observed, of the order of 1 pptv, although the vertical distribution within the boundary layer cannot be determined. Future,
simultaneous observations of O3 , HO2 and halogen oxides
from 0–1 km from an aircraft platform, therefore, would be
extremely beneficial.
Methane is removed from the atmosphere through oxidation by OH (Levy, 1971) with ∼25% of this removal occurring in the tropical MBL (Bloss et al., 2005a). In this
work the important parameters which control the OH concentration in this region have been quantified, a number of
which, until recently, have been overlooked, for example
chemical processes involving halogen species. Although the
vertical extent in the MBL of the impact of halogen oxides
on the HOx and O3 budgets, particularly for IO, remains
uncertain, these species constitute a significant source of
OH. Photolysis of HOBr and HOI lead to an increase in the
OH concentration of ∼12%, although this increase is offset slightly (∼3%) by the perturbation of the NOx chemistry (lowering [NO]) by halogen oxides as discussed in
Sect. 3.4.2 Overall, this work suggests that the methane lifetime is reduced by ∼9%, due to the presence of halogen
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L. K. Whalley et al.: Chemistry of OH and HO2 radicals over the tropical Atlantic Ocean
oxides leading to an increase in [OH]. In the current atmosphere, one added methane molecule absorbs infra-red radiation ∼25 times more efficiently than one additional CO2
molecule (Lelieveld et al., 1993) as CO2 is approximately
200 times more abundant in the atmosphere and, as a consequence, many of its absorption bands are saturated. Current
estimates suggest that methane contributes a radiative forcing of +0.48 (±0.05) Wm−2 (Forster et al., 2007), which is
∼33% of that due to CO2 . Radiative forcing from methane
increases approximately as the square root of its concentration (Forster et al., 2007), and it is estimated that without
halogen oxides, the lower [OH] would lead to an increase in
the radiative forcing by methane by ∼3.5×10−3 Wm−2 . A
key remaining question is the range of altitudes over which
halogen oxides, particularly IO, extend. To fully determine
the global impact of the halogen oxides on the oxidising capacity of the tropical BL, manifested through their effect on
[OH], the vertical profile of these species in this region is
required.
5
Summary
Measurements of OH and HO2 radicals, together with a suite
of supporting measurements, were made at the Cape Verde
Atmospheric Observatory on the island of Sao Vicente during May and June 2007.
A model using the Master Chemical Mechanism, and constrained to the supporting measurements, was used to predict OH and HO2 concentrations and to probe the chemistry
controlling the radicals at the observatory. The halogen oxides, despite their low concentrations, were found to act as
a significant HO2 sink (19%), with the subsequent photolysis of the hypohalous acids increasing the modelled [OH].
The model was able to capture daytime radical concentration
reasonably well, with the model falling within the 1 σ measurement uncertainty of 20%. During the night, the model
significantly under-predicts [HO2 ], we suggest that the thermolysis of 100 pptv of PAN, entrained from the cooler air
aloft could account for the model-measurement discrepancy
at this time.
Active halogen chemistry over the tropical oceans, where
photochemical oxidation is most prevalent, can significantly
impact the global oxidising capacity via its effect on [OH].
From this work, the calculated methane lifetime could be as
much as 9% shorter in this region if halogen chemistry is included in the calculation of [OH]. Currently, halogen chemistry is not routinely included in chemistry-transport models
used to predict, for example, global distributions of tropopsheric methane and O3 .
Further measurements of OH and HO2 coupled with concurrent measurements of halogen oxides, VOCs, HCHO,
HONO, NOx , NOy and aerosol number and size distributions are needed to describe the chemistry of this region more
fully and further test the hypotheses raised here. Vertical
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1573
profiles of the concentrations of IO, BrO, OH, HO2 and O3
are needed to discover if halogen chemistry has an impact
on the budgets of OH, HO2 and O3 throughout the boundary
layer.
Acknowledgements. The authors would like to thank P. HalfordMaw and staff of the mechanical and electronic workshops within
the School of Chemistry at the University of Leeds for their technical assistance in maintaining and developing the FAGE instrument.
Thanks also to G. McFiggans for leading the RHaMBLe project and
L. Mendes and B. Faria for campaign and logistical organisation
and to C. Seal for assistance with pack-up. The NERC SOLAS
project is acknowledged for the funding of the CVAO and NCAS
for the long-term support of the FGAM FAGE instrument. This
work was supported by NERC under grant number NE/D006589/1.
Edited by: R. von Glasow
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